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Vol. 6 (1.1) 2014

G

G e o l o g i c a l

F i e l d

T r i p s

Società Geologica Italiana

SERVIZIO GEOLOGICO D’ITALIA

Organo Cartografico dello Stato (legge N°68 del 2-2-1960)

Dipartimento Difesa del Suolo Istituto Superiore per la Protezione e la Ricerca Ambientale

ISPRA

ISSN: 2038-4947

From passive margins to orogens: the link between Zones of Exhumed Subcontinental Mantle and (U)HP metamorphism

10^thInternational Eclogite Conference - Courmayeur (Aosta, Italy), 2013

DOI: 10.3301/GFT.2014.01

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publishing group

GFT - G e o l o g i c a l F i e l d Tr i p s

Editorial Board

M. Balini, G. Barrocu, C. Bartolini, D. Bernoulli, F. Calamita, B. Capaccioni, W. Cavazza, F.L. Chiocci,

R. Compagnoni, D. Cosentino,

S. Critelli, G.V. Dal Piaz, C. D'Ambrogi, P. Di Stefano, C. Doglioni, E. Erba, R. Fantoni, P. Gianolla, L. Guerrieri, M. Mellini, S. Milli, M. Pantaloni, V. Pascucci, L. Passeri, A. Peccerillo, L. Pomar, P. Ronchi, B.C. Schreiber, L. Simone, I. Spalla, L.H. Tanner, C. Venturini, G. Zuffa.

geological field trips 2014 - 6(1.1)

Periodico semestrale del Servizio Geologico d'Italia - ISPRA e della Società Geologica Italiana Geol.F.Trips, Vol.6 No.1.1 (2014), 61 pp., 26 figs. (DOI 10.3301/GFT.2014.01)

From passive margins to orogens: the link between Zones of Exhumed Subcontinental Mantle and (U)HP metamorphism

10^th International Eclogite Conference, Courmayeur (Aosta, Italy) - Pre-conference excursions: September 2-3, 2013 Marco BELTRANDO⁽¹⁾, Roberto COMPAGNONI⁽¹⁾, Jaime BARNES⁽²⁾, Maria Luce FREZZOTTI⁽³⁾, Daniele REGIS⁽⁴⁾,

Gianluca FRASCA⁽⁵⁾, Marnie FORSTER⁽⁶⁾, Gordon LISTER⁽⁶⁾

(1) Department of Earth Sciences, University of Torino, via Valperga Caluso 35, 10125 Turin, Italy

(2) Department of Geological Sciences, Jackson School of Geosciences, University of Texas at Austin, 2275 Speedway Stop C9000, Austin, Texas 78712 (USA)

(3) Department of Earth and Environmental Sciences, University of Milano Bicocca, P.zza della Scienza 4, 20126 Milan, Italy

(4) Department of Environment, Earth and Ecosystems, The Open University, Walton Hall, Milton Keynes MK7 6AA, United Kingdom

(5) Géosciences Rennes, UMR 6118, Université de Rennes 1, Campus de Beaulieu, 35042 Rennes Cedex, France

(6) Research School of Earth Sciences, Mills Rd, Australian National University, Canberra, ACT 0200, Australia Corresponding Author e-mail address: marco.beltrando@unito.it

Responsible Director

Claudio Campobasso (ISPRA-Roma) Editor in Chief

Gloria Ciarapica (SGI-Perugia) Editorial Responsible

Maria Letizia Pampaloni (ISPRA-Roma) Technical Editor

Mauro Roma (ISPRA-Roma) Editorial Manager

Maria Luisa Vatovec (ISPRA-Roma) Convention Responsible

Anna Rosa Scalise (ISPRA-Roma) Alessandro Zuccari (SGI-Roma)

ISSN: 2038-4947 [online]

http://www.isprambiente.gov.it/it/pubblicazioni/periodici-tecnici/geological-field-trips

The Geological Survey of Italy, the Società Geologica Italiana and the Editorial group are not responsible for the ideas, opinions and contents of the guides published; the Authors of each paper are responsible for the ideas, opinions and contents published.

Il Servizio Geologico d’Italia, la Società Geologica Italiana e il Gruppo editoriale non sono responsabili delle opinioni espresse e delle affermazioni pubblicate nella guida; l’Autore/i è/sono il/i solo/i responsabile/i.

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INDEX

Information

Safety ...5

Hospital ...5

Accomodation ...5

Riassunto ...7

Abstract ...8

Program Summary ...9

Excursion notes 1. Excursion aims ...11

2. Jurassic paleogeography of the Alpine Tethys ...13

DAY 1 3. Punta Rossa unit (Valaisan domain): Mesozoic rift-related juxtaposition of sub-continental mantle, continental basement and sediments and their interaction during Alpine metamorphism ...14

3.1 Introduction ...14

3.2 Geological Setting ...14

3.3 Lithostratigraphy of the Punta Rossa unit ...16

3.3.1 Ultramafic rocks, Paleozoic basement and Mesozoic pillow lavas ...18

3.3.2 Metasediments ...20

3.4 Hermite unit ...21

3.5 Tectono-metamorphic evolution ...22

3.5.1 High pressure metamorphism pillow lavas ...22

3.5.2 Post-high pressure tectono-metamorphic evolution .22 DAY 2 4. UHP metamorphism of early post-rift sediments in the Lago di Cignana unit, Zermatt-Saas zone, Valtournenche ...26

4.1 Geological setting ...26

4.1.1 Deformation history/style ...26

4.1.2 Main lithologies of the Lago di Cignana unit ...27

4.1.3 The LCU metamorphic evolution ...28

4.1.4 Age of UHP metamorphism ...30

4.2 The impure quartzite and the manganiferous garnetite ...30

4.2.1 Garnets of garnetite ...31

4.2.2 Inclusions in garnet of garnetite ...31

Itinerary Day 1: Punta Rossa unit (Valaisan domain) STOP 1.1 - Alpine architecture: The Punta Rossa-Hermite-PSB stack ...33

STOP 1.2 - Contact between Punta Rossa and Hermite units ..34

STOP 1.3 - A paleo-ocean floor: the contact between ultramafic rocks and metasedimentary breccia ...37

STOP 1.4 - Serpentinized spinel peridotite ...39

STOP 1.5 - The contact between continental basement and ultramafic rocks: pre-metamorphic brittle deformation and alpine metasomatism ...41

STOP 1.6 (optional) - (Overturned) stratigraphic transition between Paleozoic basement, Mesozoic breccia and radiolaria-bearing garnet-chloritoid micaschist ...43

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DAY2: Lago di Cignana unit, Zermatt-Saas zone, Valtournenche

STOP 2.1 - Ophiolites of the Zermatt-Saas/Combin zone and Etirol-Levaz continental crust slice ...46 STOP 2.2 - The UHP Lago di Cignana unit (LCU) ...47

Stop 2.2a: Eclogites exposed to the south of the lake

shore close to the dam ...47 Stop 2.2b: Metasedimentary siliceous and carbonate

rocks exposed along the lake shore ...49 Stop 2.2c: Thin layers of sheared metagabbro with

fuchsite pseudomorphs after original Cr-clinopyroxene at the upper contact with the Zermatt-Saas zone ...52 References ...56

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Information Day 1

Altitude: 2100-2550 m a.s.l.

Elevation gain (on foot): ca. 150 m along an easy path, followed by ca. 300 m on grassy slopes

Day 2

Altitude: ca. 2150 m a.s.l.

Elevation gain (on foot): none Safety

Mountain boots, sun glasses and sunscreen are strongly recommended.

Considering that sudden weather changes are likely to happen in Alpine settings, water and windproof clothes are also recommended.

Hospital Aosta

Ospedale Regionale Viale Ginevra 3, Aosta, Tel. +39 0165 5431 ý 0165 5431

Accomodation Hotel Pavillon

Strada Regionale 62, 11013 Courmayeur Tel: +39 0165 846120

Location of the two excursion days

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DAY 1

DAY 2

Localization of the route of the first day Localization of the route of the second day

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Riassunto

In questa escursione verranno trattati alcuni temi connessi all’evoluzione della crosta ofiolitica delle Alpi Occidentali dal rifting Giurassico al successivo metamorfismo di (ultra-)alta pressione nell’Eocene-Oligocene.

Le tematiche principali verteranno (1) sull’architettura delle unità di (U)HP, (2) sull’evoluzione deformativa polifasica durante la subduzione/esumazione e (3) sul significato geodinamico circa la presenza di scaglie di basamento continentale giustapposte a peridotiti serpentinizzate ed a sedimenti di sin- e post-rift, comune a molte unità ofiolitiche alpine di alta pressione. Nel corso della prima giornata, al Passo del Piccolo San Bernardo, osserveremo le relazioni tra basamento continentale Paleozoico, mantello sotto-continentale serpentinizzato e sedimenti di sin- e post-rift, stabilitesi prima dell’orogenesi alpina. Tali relazioni sono preservate nonostante un’evoluzione deformativa polifasica accompagnata da metamorfismo a P≥1.5GPa.

Relazioni analoghe tra questi litotipi sono tipiche delle “Zones of Exhumed Subcontinental Mantle” (ZESM) presenti lungo i margini attuali poveri in magma. In tali contesti il mantello sotto continentale viene esumato per l’attività di sistemi di faglie a basso angolo e grande rigetto. Tale tettonica estensionale può portare al campionamento di sezioni di basamento continentale di tetto, dando così origine ai cosidetti ‘alloctoni estensionali’. I ritrovamenti effettuati in alcune unità alpine ad ofioliti, quindi, indicano che associazioni litostratigrafiche che consistono di mantello sotto-continentale serpentinizzato, basamento continentale e sedimenti di sin- e post-rift non sono necessariamente legate a dinamiche subduttive complesse, ma possono altresì essere ereditate dalle fasi estensionali pre-subduzione. Nel corso della seconda giornata tratteremo l’evoluzione tettono-metamorfica di sedimenti di post-rift, basalti di fondo oceanico nell’unità di (ultra-)alta pressione del Lago di Cignana, con particolare attenzione ai recenti ritrovamenti di micro-diamanti inclusi nei granati di noduli in quarziti manganesifere. In quest’area la deformazione polifasica durante gli stadi esumativi è stata accomodata da varie generazioni di zone di taglio, accompagnate dalla formazione di pieghe. Per questa ragione non è stato fino ad ora possibile riconoscere sezioni coerenti di crosta giurassica analoghe a quella mostrata nel corso del primo giorno di questa escursione.

Parole chiave: Alpi Occidentali, margini di rift poveri in magma, eclogite, eredità strutturale, quarziti a granatiti manganesifere, diamante.

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Abstract

This excursion deals with the evolution of the ophiolite-bearing crust of the Western Alps from Jurassic rifting to Eocene-Oligocene (ultra-)high-pressure metamorphism. The main topics of this excursion are: (1) the architecture of (U)HP units; (2) multi-stage deformation during subduction/exhumation; (3) the geodynamic significance of continental basement juxtaposition with serpentinized peridotites and syn- to post-rift sediments, which is common to several alpine (U)HP ophiolites.

During the first day, at the Petit St. Bernard pass, we will show that continental basement, serpentinized subcontinental mantle and metasediments preserve pre-Alpine rift-related relationships, despite having undergone multiple phases of alpine deformation and metamorphism at P≥1.5GPa. Similar relationships are typical of Zones of Exhumed Subcontinental Mantle (ZESM), observed along present-day magma-poor rifted margins, where sub-continental mantle is exhumed at the basin floor through the activity of low-angle detachment faults. Fault activity may locally result in sampling of hangingwall material, including continental basement, which then resides on the fault plane as an “extensional allochthon”. These findings indicate that the presence of serpentinized ultramafics alongside continental basement and syn- to post-rift sediments, which is common to several alpine metamorphic terrains, is not necessarily related to complex subduction dynamics, but it may also be inherited from the pre-subduction history.

During the second day, a visit to the Lago di Cignana ultra-high pressure unit will give us the opportunity to discuss the tectono-metamorphic evolution of early-post rift metasediments and metabasalts in the light of the recent finding of microdiamond inclusions in garnet from Mn-bearing quartzites. Multi-stage exhumation- related shuffling of the tectono-metamorphic pile, together with late-stage folding, prevents the recognition of coherent sections of Jurassic crust of the kind observed during the first day.

Keywords: Western Alps, magma-poor rifted margin, eclogite, structural inheritance, quartzite with manganesiferous garnetite, diamond.

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Program summary

The Punta Rossa unit, which will be visited on the first day, crops out in the Breuil Valley, near the Petit St. Bernard Pass, at the French-Italian border (Fig. 1). The pass will be approached from La Thuile, in Italy.

The western side of the Breuil Valley, which is the target of this excursion, can be reached following track number 14, which starts at the parking area near Lac de Verney, close to the National Route 26.

To reach this parking area, driving up from La Thuile, leave the main road at a sharp left bend at the altitude of 2110 m, when you are already in sight of the pass and drive for 50 meters on an unsealed road heading right. Due to the relatively high altitude of the area, in the 2100-2550 m range, this fieldtrip can only be

Fig. 1 - Topographic map of the Petit St. Bernard Pass area with access route

and excursion Stops.

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undertaken during the summer. Please, note that many outcrops can still be covered in snow in June. August to early September is normally the best time of the year to avoid any snow cover. Several outcrops described in this field guide are located away from the main track and can be accessed only by walking on moderately steep grassy and rocky slopes. Therefore, adequate hiking gear is recommended, as well as warm and water proof clothes, as this area is often windy and weather changes can occur rapidly. Mobile phone coverage in the Breuil valley is generally good. Also note that the area is devoid of suitable shelters in case of unexpected storms.

To reach Lago di Cignana on the second day, drive to Antey Saint André, in Valtournenche, then take the road to Torgnon up to the locality Champtornè-Dès (1885 m a.s.l.).

Only minivans and cars can continue from here along an unsealed road up to the Lago di Cignana dam (Fig. 2). Please note that a permit must be requested to the Torgnon municipality to drive on this road. This drive is about 15 km long. All excursion stops are located in close vicinity to the road.

Fig. 2 - Topographic map of the Lago di Cignana area with access route and excursion Stops.

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excursion notes

1. Excursion aims

Orogenic belts, sampling fossil subduction zones, provide key natural laboratories to investigate convergent margin dynamics (e.g. Agard et al., 2009; Beltrando et al., 2010a). In this context, elucidating the pre- orogenic lithostratigraphy of the tectono-metamorphic units stacked in the mountain belt has important implications for understanding the processes leading to their sampling from the downgoing plate and their subsequent exhumation.

In the high-pressure part of Alpine-type orogenic belts, ophiolite-bearing units often consist of Paleozoic continental basement slivers cropping out alongside Jurassic ophiolites and syn- to post-rift sediments (e.g.

Dal Piaz, 1999; Beltrando et al., 2010b and 2012; Meresse et al., 2012; Vitale Brovarone et al., 2011). These seemingly ‘anomalous’ lithological associations are commonly ascribed to complex orogen dynamics, including centrifuge-like flow in a subduction channel (e.g. Polino et al., 1990; Gerya et al., 2002). However, studies conducted over the past 30 years both in present-day rifted margins (e.g. Peron Pinvidic & Manatschal, 2009) and in fossil analogues preserved in the weakly deformed Eastern Alps (Froitzheim & Manatschal, 1996;

Manatschal, 2004; Manatschal et al., 2006) provided compelling evidence of the existence of a specific type of transitional lithosphere between typical ‘oceanic’ and ‘continental’ lithosphere. These transitional domains, which can be up to 200 km wide, comprise the distal continental margin and the so-called “Zone of Exhumed Subcontinental Mantle” (ZESM). In the distal continental margin, thinned continental crust (<10 km) is directly onlapped by post-rift sediments. As continental crust thins out oceanward, this area grades into the ZESM, where subcontinental mantle is exhumed at the seafloor by the activity of long-offset detachment faults. As a result of the activity of detachment systems, serpentinized mantle can be overlain by slivers of continental crust derived from the stretched hangingwall. The resulting basement topography exerts a strong control on the distribution of syn-rift sediments, which are then sealed by typical pelagic post-rift sediments. Typical distal continental margins and ZESM’s have already been extensively recognized in sections of peri-Mediterranean mountain belts that were only marginally affected by orogeny-related deformation/metamorphism (e.g.

Decandia & Elter, 1972; Froitzheim & Eberli, 1990; Florineth & Froitzheim, 1994; Molli, 1996; Durand-Delga et al., 1997; Marroni et al., 1998; Manatschal, 2004; Mohn et al., 2012; Beltrando et al., 2013). More recently, lithological associations with slivers of continental basement resting upon serpentinized ultramafics in the high- pressure to ultra-high-pressure Piemonte units of the Western Alps and Corsica have been interpreted as

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excursion notes

remnants of Jurassic ZESM’s (Dal Piaz, 1999;

Beltrando et al., 2010b; Vitale Brovarone et al., 2011; Meresse et al., 2012). Similar findings in the Valaisan domain of the Western Alps (Loprieno et al., 2011; Beltrando et al., 2012) suggest that ZESM’s may be much more common than hitherto recognized.

During excursion days (Fig. 3) we will first visit the best-preserved example of a Mesozoic ZESM sampled within the Western Alps, at the Petit St. Bernard Pass. In this locality, a combination of structural, petrologic and lithostratigraphic observations indicates that the different rock types were already juxtaposed prior to the Alpine orogeny. In the Lago di Cignana area, instead, Alpine deformation largely erased pre-orogenic relationships among the different rocks types.

However, several lines of evidence indicate that this (U)HP unit was originally part of a

“Zone of Exhumed Subcontinental Mantle”, too.

Fig. 3 – Simplified tectonic map of the Western Alps (from Beltrando et al., 2010). BI: UHP Brossasco-Isasca unit; C: Combin unit; GP: Gran Paradiso Massif; MB: Mont Blanc Massif; MR: Monte Rosa Massif; V: Monviso; Z: Zermatt-Saas unit; ZH:

Houiller Zone. The areas that will be visited during excursion days are indicated.

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excursion notes

2. Jurassic paleogeography of the Alpine Tethys

Since the early ‘80’s, a large number of studies has been devoted to constraining the pre-orogenic paleogeography of the area that was later involved into the Alpine orogeny in the Western Mediterranean (see Mohn et al., 2010, Masini et al., 2013 and references therein). The generally accepted view of the paleogeography of this domain consists of two continental masses, namely the Adriatic/African plate and the European plate, separated by an oceanic basin, labeled Western Tethys (Fig. 4). This basin was characterized by a main branch, known as South Penninic (or Piemonte-Liguria) basin, and a minor northerly branch, the North Penninic (or Valaisan) basin (e.g. Trümpy, 1980; Lemoine, 1985). The two sub-basins, in the area that later became part of the Western and Central Alps, were separated by a continental high, the Briançonnais rise (e.g. Trümpy, 1949; Lemoine, 1985). Several paleogeographic reconstructions suggest that the Western Tethyan basin was highly fragmented by regional-scale EW-trending transform faults (e.g. Lemoine, 1985).

Subsequent studies performed in little deformed parts of the Alpine belt have progressively refined this early picture (Fig. 4), showing that large areas of the Piemonte-Liguria and Valaisan basins were floored by partly serpentinized sub-continental mantle (e.g. Manatschal & Müntener, 2008) and that thinned continental basement directly overlain by syn- to post-rift sediments marked the transition from the proximal Adriatic margin to the Piemonte-Liguria basin (Froitzheim & Eberli, 1990; Masini et al., 2011). The resulting paleogeography, chiefly based on observations performed in the little deformed Eastern Alps, consists of an Adriatic proximal margin, preserved in the upper Austroalpine nappes, grading

outboard into a necking zone, preserved in the middle Austroalpine nappes (Mohn et al., 2012) then

followed by the distal continental margins, floored by exhumed crust, now sampled in the lower Austroalpine nappes. The areas floored by exhumed mantle belonging to the South Penninic basin are now best preserved in the Forno, Platta and Totalp units.

Fig. 4 –

Western Tethys paleogeography (modified after Mohn et al., 2010). Note that the Adriatic and

European margins, which are currently stacked in the Western Alps, originally were not conjugate.

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DAY 1

3. Punta Rossa unit (Valaisan domain): Mesozoic rift-related juxtaposition of sub-continental mantle, continental basement and sediments and their interaction during Alpine metamorphism 3.1 Introduction

The Valaisan domain, near the Petit St. Bernard Pass, consists of serpentinized sub-continental mantle juxtaposed to Paleozoic basement, Mesozoic pillow meta-basalts and Mesozoic to Tertiary meta-sediments, which underwent Alpine metamorphism at P≥1.5 GPa (Cannic et al., 1996; Bousquet et al., 2002). A recent structural, lithostratigraphic and metamorphic study (Beltrando et al., 2012) showed that this complex lithostratigraphic association was largely established during rift-related extensional tectonics and underwent relatively minor reworking during the Alpine orogeny. Mantle peridotites were exhumed at the bottom of the North Penninic basin by extensional faulting, resulting in widespread cataclasis of continental basement rocks, which rested above serpentinized mantle as extensional allochthons. The serpentinite-Paleozoic basement pair was sealed by locally-sourced polymictic breccias, prior to the deposition of radiolarian-bearing gray micaschists, followed by other basinal metasediments, including calcschists. Despite subsequent Alpine deformation and metamorphism, resulting in multi-stage folding and medium P-low T metamorphism, the rift- related relationships among the different rock types can still be observed or inferred in several localities. This field guide presents a one day excursion aimed at illustrating some of the key features that allow reconstructing the pre-Alpine lithostratigraphy, alongside a few stops where the present-day architecture of this domain is described. The description of the serpentinites is also enriched by geochemical data, which will be presented during the excursion, documenting the origin of these mantle peridotites as well as the multiphase hydration during subduction and metamorphism, through interaction with sediment-derived fluids.

3.2 Geological setting

The Valaisan domain, in the Western Alps, is bounded towards the NW by the Penninic Front, which separates the Valaisan units from the more external Helvetic-Dauphinois domain (Fig. 3). The latter consists of Paleozoic basement overlain by a Late Permian to Lower Oligocene sedimentary cover sequence, typical of the European

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proximal rifted margin (Escher et al., 1997). Towards the SE, instead, the Houiller Front separates the Valaisan units from the overlying Houiller Zone (Fig. 3 and 5). The latter belongs to the Briançonnais domain, which consists of tectonic units originated from a continental rise separating the North and South Penninic basins. The Houiller Zone consists of Carboniferous and Permian sandstone, breccia and shale of continental origin, with minor Early Triassic dolomite and shale, which underwent relatively minor Alpine metamorphism at P= 0.6 ± 0.2 GPa and T= 280-300°C (Lanari et al., 2012).

Fig. 5 – Simplified geological map of the Valaisan units in the Breuil Valley. Inset shows the different tectono-stratigraphic units detected in the area (from Beltrando

et al., 2012).

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excursion notes

The Valaisan units, in the Petit St. Bernard pass area, at the French-Italian border, have been the subject of several recent studies (Cannic et al., 1995 and 1996; Fügenschuh et al., 1999; Schärer et al., 2000; Bousquet at al., 2002;

Beltrando et al., 2007; Masson et al., 2008; Mugnier et al., 2008; Loprieno et al., 2011; Beltrando et al., 2012 and references therein). The pre-Alpine lithostratigraphy of the Valaisan domain between the Penninic front and the Petit St. Bernard pass indicates an original paleogeographic position along the European escarpment, between the European plate sensu stricto and the Briançonnais rise (Fig. 4; Loprieno et al., 2011). Evidence of mantle exhumation at the seafloor indicates that complete crustal excision was achieved, at least locally, within this basin (Loprieno et al., 2011; Beltrando et al., 2012). The Valaisan domain, in the Petit St. Bernard area, can be subdivided into the Hermite and Punta Rossa sub-units (Figs. 5, 6 and 7), which are separated by a late-stage shear zone (Stop 1.2) (Beltrando et al., 2012). Despite this subdivision, the two units display a similar tectono- metamorphic evolution and lithostratigraphy. However, serpentinites are exclusively found in the Punta Rossa unit, which is the main focus of this excursion.

In the Petit St. Bernard pass area, the small Petit St. Bernard unit crops out between the Houiller zone and the Valaisan domain (Fig. 5). The Petit St. Bernard unit consists of marbles and dolostones, originated from upper Triassic dolomitic schists and Lower Jurassic cherty marbles, overlain by calcschists (Elter & Elter, 1957). The calcschists consist of belemnite-bearing carbonate micaschists and graphitic marbles (Franchi, 1899), interpreted as Early to Middle Jurassic in age (Elter & Elter, 1965). This unit has generally been considered as an independent Alpine tectono-metamorphic unit, separated from the mafic/ultramafic bearing units underneath by Alpine shear zones (e.g. Elter & Elter, 1957, Antoine et al., 1992; Bousquet et al., 2002, Beltrando et al., 2012; see Masson et al., 2008 and Loprieno et al., 2011 for alternative interpretations). This view is supported by the distinctive lithostratigraphy and by its tectono-metamorphic evolution (Beltrando et al., 2012), including the estimated timing of Alpine metamorphism, in the 43-45 Ma range (⁴⁰Ar-³⁹Ar step heating on white mica; Cannic, 1996).

3.3 Lithostratigraphy of the Punta Rossa unit

The Punta Rossa unit consists of serpentinized subcontinental mantle, Paleozoic basement, rare Mesozoic pillowed meta-basalts and different types of Mesozoic metasediments. Only the general features of the different rock types are recalled here, as a more detailed discussion is presented in the excursion stops.

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excursion notes

Fig. 6 - Geological map with excursion stops (from Beltrando et al., 2012). Traces of geological cross sections shown in Figure 7 are also indicated.

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3.3.1 Ultramafic rocks, Paleozoic basement and Mesozoic pillow lavas

Serpentinized ultramafic rocks crop out throughout the Breuil Valley. The largest outcrops of serpentinized ultramafics are located near Punta Rossa (Stops 3 to 5) and further to the north, at Tormottaz Lake (Figs. 7 and 8). Serpentinite bodies display textural and mineralogical zoning, dependent upon the distance from the surrounding polymictic breccias and continental basement rocks.

Fig. 7 - Geological cross sections with excursion stops. Location of cross sections is indicated in Figure 6. Grey arrows indicate stratigraphic polarity. The pre-Alpine geometry, as inferred from structural and lithostratigraphic observations, is also depicted (from Beltrando et al.,

2012).

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The significance of this zoning in terms of Jurassic and Tertiary evolution is discussed in Stops 1.3 and 1.4.

Laterally discontinuous slivers of crystalline basement rocks are commonly associated with the ultramafic rocks (Figs. 5-8). Crystalline basement largely consists of leucocratic peraluminous metagranitoids (Fig. 9a), whose igneous protolith has been dated at 267±1 Ma (Beltrando et al., 2007). Paleozoic basement also consists of mafic–ultramafic sills and laccoliths interbedded with black schists, which are referred to as ‘Versoyen complex’

(Fig. 9b; Elter & Elter, 1965). Sills range in thickness from 0.5 to 40 m and locally show internal zoning, with cumulate gabbros and ultramafic layers passing upward to dolerites (Cannic, 1996). Chilled margins and intrusion breccias are common at the contact with the black schists (Beltrando et al., 2007). The cores of sills have flat REE patterns characteristic of N-MORB and T-MORB, while the geochemistry of the sill margins provides evidence for contamination caused by the intrusion of hot mafic magmas into unconsolidated sediments rich in water (Cannic, 1996 and Mugnier et al., 2008). As pointed out by Loubat (1975) and Kelts (1981), the Versoyen complex resembles sill–sediment complexes presently found in the Gulf of California, where they are related to high sedimentation rates during spreading within narrow oceanic pull-apart basins. Locally, metagranitoid dykes

Fig. 8 - Serpentinized ultramafics, Paleozoic continental basement and metasediments in the

Tormottaz Lake area (Beltrando et al., 2012).

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preserve intrusive relationships with black micaschists typical of the Versoyen complex, indicating that the latter is also Paleozoic (Fig. 9a).

Similar cross-cutting relationships are found in the neighboring Hermite unit (Schärer et al., 2000), where the Paleozoic age of the mafic magmas is also confirmed by U-Pb geochronology. Mafic magmatism was probably multi-stage, as the intrusion of a rather large mafic complex in the Aiguille de Clapet area (France) has been constrained at 337±4.1 Ma (Masson et al., 2008), while younger leucogabbro dykes were emplaced at 272±2 Ma (Beltrando et al., 2007).

Meta-pillowed lavas have been described from several localities within the Punta Rossa unit. Based on our observations, the only unambiguous evidence for basaltic lavas erupted at the seafloor is preserved near the Bassa Serra pass (Fig. 9c; Loubat, 1968; see Beltrando et al., 2012, for a discussion).

3.3.2 Metasediments

Serpentinized ultramafic rocks, Paleozoic basement and meta-pillow lavas are invariably associated with metabreccias, which contain clasts of all the lithologies described above. The lithological composition of this breccia is variable and is strongly controlled by the type of associated basement. Our structural study reveals that the meta-breccia outcrops found in the area were originally located along the same horizon (Fig. 7; Beltrando et al., 2012). The thickness of the metabreccia is variable, probably as a combined effect of pre-Alpine basin geometry and Alpine deformation. It ranges from a few meters, on the western side of Tormottaz Lake, where ultramafics and gray schists are separated by less than 5 m of polymictic breccia and meta-arkose, to about 20 m, on the ridge to the south-west of

Fig. 9 - Paleozoic basement consists mainly of meta-leucogranites (a), locally intrusive into black micaschists. The black micaschists are generally associated with

mafic sills (b). Mesozoic pillowed metabasalts (c) are found near Bassa Serra Col.

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Punta Rossa. In the Punta Rossa area, where continental basement is rather homogeneous, differentiating between cataclastically deformed basement and re-sedimented monogenic breccia is often very difficult, due to the extensive Alpine deformation and metamorphism. As a result, we opted for grouping together tectonic and sedimentary monogenic breccias in the map presented in Fig. 6 under the ‘tectono-sedimentary breccia’ label. In all circumstances, the first unambiguous occurrence of sedimentary breccias is marked by the presence of polygenic breccias with a black matrix. Metabreccias often preserve transitional contacts with the underlying crystalline rocks and the overlying gray micaschists (Stop 1.6). Importantly, the “Collet des Rousses metabreccia”, which is part of the Hermite unit and displays significant lithological similarities and identical stratigraphic position with respect to the metabreccias of the Punta Rossa unit, contains belemnites in the matrix, evidence for a Mesozoic depositional age.

The metabreccia is generally directly in contact with grey micaschists, consisting of graphitic schists that become progressively more carbonate-rich towards the (stratigraphically) overlying calcschists. Clasts of metagranitoids, metamafics and marbles are locally observed. The contact between carbonate-rich micaschists and the overlying calcschists is either gradational or characterized by interlayered calcschists.

A large part of the central and northern part of the Breuil Valley then consists of metasediments normally ascribed to the ‘Valaisan trilogy’ (Burri, 1979; Trümpy, 1951; see Masson et al., 2008 for a different view on the significance of these metasediments). From the contact with the gray micaschists, these metasediments are grouped in ‘Aroley marble’, ‘Marmontains quartzite’ and ‘St. Christophe calcschists’. The Aroley marble consists of impure marble and carbonate schists, with rare conglomeratic beds near the base. The

‘Marmontains quartzite’ is characterized by alternating beds of carbonate-free black schists and quartzarenites, while the ‘St. Christophe calcschists’ consist of calcareous–arenaceous strata and black marls and schists. The onset of the Valaisan trilogy sedimentation is poorly constrained, due to the rarity of fossils, and it has been placed at different stages in the Cretaceous, either in the Barremian–Aptian (Elter, 1954; Sodero, 1968) or in the Turonian–Santonian (Antoine, 1965, 1971; see Loprieno et al., 2011 for a discussion).

3.4 Hermite unit

The Hermite unit consists of Versoyen complex rocks (e.g. Loubat, 1968), Collet des Rousses metabreccias and Arguerey calcschists (Antoine, 1971). In the Aiguilles de l’Hermite-Punta dei Ghiacciai area, aegirine and glaucophane rims are observed around magmatic clinopyroxene. The Collet des Rousses metabreccia is

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characterized by a laterally variable thickness and is locally absent at the Versoyen complex–Arguerey calcschist interface. The Arguerey calcschists consist of chlorite-white mica graphitic marble, phengitic micaschist, graphitic-carbonate paragreenstone, graphitic marble interlayered with chloritic marble and carbonate schist. Locally, the Arguerey calcschists contain rounded boulders of gray marble, which range in diameter from 30 cm to 1 m. For more details on the Hermite unit see Beltrando et al. (2012).

3.5 Tectono-metamorphic evolution 3.5.1 High pressure metamorphism

High pressure mineral assemblages are only rarely preserved in the Punta Rossa and Hermite units (Fig. 10): jadeite has been described in the Punta Rossa metagranitoids (Saliot 1979), carpholite is reported from metasediments of the Petit St. Bernard and Punta Rossa units (Bousquet et al., 2002; Goffé & Bousquet, 1997; Loprieno et al., 2011) and eclogites are locally preserved in large layered mafic complexes, which largely escaped exhumation-related deformation (Schürch, 1987; Cannic et al., 1996). Static re-equilibration of Paleozoic Fe-Ti-rich gabbros in the Hermite unit resulted into pseudomorphic replacement of omphacite (Jd50-55) after augite, Ca-Fe rich garnet after plagioclase and rutile after ilmenite, variably associated with glaucophane 1 ± zoisite/clinozoisite ± quartz (Fig. 10;

Cannic et al., 1996). Fe-Mg partitioning between co-existing garnet and clinopyroxene, which are never found directly in contact, combined with the phengitic substitution in rare white mica, yielded metamorphic conditions of P=1.4-1.6 GPa at T=425-475°C (Cannic et al., 1996). This temperature estimate lies at the upper limit of the maximum temperatures for metamorphism indicated by graphite crystallinity in Mesozoic and Paleozoic metasediments, falling in the 390-430°C range (Beltrando et al., 2012). Black micaschists of the Versoyen complex and grey micaschists of the Mesozoic cover are generally strongly re-equilibrated at lower-P conditions (see below), but rare lawsonite relicts have been observed (Cannic et al., 1996). In metasediments the carpholite + phengite + chlorite + quartz assemblage found in metamorphic veins has been estimated to have formed at P=1.5-1.7 GPa and T=350-400°C (Bousquet et al., 2002).

3.5.2 Post-high pressure tectono-metamorphic evolution

Post-high pressure metamorphism, in the metagabbros, is recorded by crystallization of a glaucophane 2 + zoisite/clinozoisite + phengite + paragonite + albite mineral assemblage, followed by re-equilibration at greenschist facies conditions, as indicated by the common assemblage actinolite-tremolite + albite + chlorite + Fe-epidote ±

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stilpnomelane ± pumpellyite (Cannic et al., 1996). Static greenschist facies retrogression is very pervasive in the relatively small mafic sills that will be observed during this excursion. Due to Alpine metamorphism, the mafic layers of the Versoyen complex now consist of albite–chlorite–actinolite bearing metamicrogabbros, albite–chlorite–epidote bearing metadolerites and greenstones, which only rarely preserve evidence of pre-existing glaucophane. As a result, the post-high pressure tectono-metamorphic evolution is best observed in the radiolaria-bearing grey micaschists, which preserve micro- and mesoscale evidence of the multistage post-HP deformation and metamorphic history (Fig. 11). In this lithology, relics of the HP evolution are restricted to (1) lozenge-shaped aggregates of muscovite, paragonite and chlorite, possibly derived from earlier lawsonite, (2) prismatic aggregates of chloritoid, quartz, white mica and chlorite, which may derive from former carpholite (Fig. 11a; Bousquet et al., 2002) and (3) a relict foliation, preserved in microlithons, defined by white mica with a high celadonite content (Si=3.4–3.5 a.p.f.u.). These relics are largely overprinted by a pervasive planar fabric, defined by white mica (Si=3.3 a.p.f.u.) and chloritoid, which is axial planar to isoclinal folds that affect the entire lithostratigraphy (Fctd in Beltrando et al., 2012 and F1 in Fügenschuh et al., 1999 and Loprieno et al., 2011). The ubiquitous preservation of the original lawsonite and chloritoid porphyroclast shape is interpreted as evidence that their re-equilibration post-dated the formation of Sctd. A second generation

Fig. 10 - High Pressure mineral assemblages are rarely preserved. Carpholite is locally found with quartz in metamorphic veins (a), while omphacite relicts have been reported exclusively from the Clapet gabbro, in the Hermite unit (b, from Cannic et al., 1996). The PT estimates proposed for metamafics (Cannic et al., 1996) and for metapelites (Bousquet et al., 2002) are indicated by red and violet rectangles, respectively (c). The minimum pressure of metamorphic re-equilibration of the metaleucogranite is constrained by the jadeite-in reaction. The maximum T of

metamorphism is constrained by Raman thermometry (Beltrando et al., 2012).

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of chloritoid crystals is oriented at a high angle with respect to Sctd, often associated to small garnet porphyroblasts, with a diameter <100 μm (Fig. 11a). P=1.3-1.5 GPa and T= 450-500°C have been proposed for the replacement of Fe-carpholite by chloritoid + chlorite + phengite (Bousquet et al., 2002). However, as noted above, the estimated temperatures greatly exceed those determined by thermometry based on Raman spectroscopy of carbonaceous material on the same metasediments (Beltrando et al., 2012). ⁴⁰Ar/³⁹Ar geochronology on metagranitoids and metasediments characterized by a pervasive Sctd consistently yielded apparent ages in the 32.8-34.2 Ma range (Cannic, 1996), indicating that HP metamorphism was pre-Oligocene in age.

Sctd is affected by mesoscale folds with rounded hinge, axial plane dipping to the west (270/40) and fold axis oriented NS (Fig. 11c). Such folding is associated with the formation of a new axial planar cleavage, with sub- mm to mm spacing, normally restricted to the hinge area. The new fabric is defined by white mica, chlorite, quartz and graphite and is labeled Srec (where ‘rec’ indicates ‘recumbent’) in Beltrando et al. (2012) and F2 in Fügenschuh et al. (1999) and Loprieno et al. (2011). Folding is also responsible for deformation of the white mica aggregates developed at the expense of porphyroclastic lawsonite.

Formation of large-scale recumbent folds is followed by localized shearing along the Punta Rossa-Hermite units contact, which is slightly discordant with respect to the pre-existing lithological boundaries and to the axial planes of Fctd and Frec. Microscopic observations (Fig. 13c) indicate that the SE dipping chlorite-rich shear planes (labeled SZSE in Fig. 13), which are common within 2–4 m from the contact, cut across chlorite-rich pseudomorphs developed at the expense of former chloritoid. These observations are taken to indicate that top-to-the-southeast shearing took place at greenschist facies conditions. Late-stage shearing is then followed by large-scale upright folds (Fup in Beltrando et al., 2013 and F3 in Loprieno et al., 2011) characterized by steep axial planes and NNE–SSW to NE–SW striking fold axes (Fig. 11d). Fup folds are characterized by a box fold geometry and by relatively large along-axis strain gradient, resulting in significant variations in the fold amplitude. At the outcrop scale, Fup folds can locally be associated with a spaced axial planar cleavage restricted to a few outcrops to the north-west of the Tormottaz peak.

The overall architecture of the Hermite and Punta Rossa units, in the Breuil Valley, is mainly controlled by two large recumbent folds (Frec/F2) deformed by an antiform-synform pair related to Fup/F3 (Fig. 7). The IEC excursion will take place on a limb of a major F2 fold, resting on the SE limb of the large F3 Tormotta anticline (Fig. 7). Therefore, in the area visited during the excursion, the effects of F2 and F3 are relatively minor and F1 folds can be observed, providing access to the oldest stages of the tectono-metamorphic evolution of the Valaisan domain.

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Fig. 11 - Tectono- metamorphic evolution recorded in grey

micaschists. In this lithology the main fabric is marked by chloritoid and

white mica and is axial planar to isoclinal folds (Sctd; a, b).

This fabric wraps around relict carpholite and lawsonite and is statically

overgrown by garnet and by a second generation of chloritoid. Sctd is then deformed by two main fold generations (c and d).

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DAY 2

4. UHP metamorphism of early post-rift sediments in the Lago di Cignana unit, Zermatt-Saas zone, Valtournenche

4.1 Geological setting

The ultrahigh pressure (UHP) Lago di Cignana unit (LCU) is exposed in the upper Valtournenche, Aosta Valley, Italian Western Alps (Figs. 3 and 18). The LCU is part of the Piemonte zone of calcschists with meta-ophiolites.

It is considered to derive from the Jurassic Tethys Ocean, which separated the European continent from the Apulia (or Adria) plate (Fig. 4; Dal Piaz, 1974; Dewey et al., 1989; Polino et al., 1990; Lombardo et al., 2002 and references therein). In the upper Valtournenche, the Piemonte zone consists of a pile of tectonic slivers including both epidote–blueschist to lawsonite–blueschist facies (Combin zone) and eclogite facies (Zermatt–Saas zone) Alpine metamorphic rocks as defined by Bearth (1967). These two main ensembles of tectonic units of meta- ophiolites and metasediments are sandwiched between two continental crust units, the overlying Austroalpine Dent Blanche and the underlying Penninic Monte Rosa Massif (Fig. 3). The LCU is best exposed on the southern side of the artificial Lago di Cignana (Figs. 18, 19). Its small size, first recognized by Reinecke (1998), was later reduced by detailed mapping by Tamagno (2000) and Forster et al. (2004) (Fig. 19). The unit, which is <100 m thick, consists of 3 main slivers ~ 1000, 350 and 250 m long (Forster et al., 2004). It is overlain by a thin unit consisting of highly deformed metagabbro belonging to the Zermatt-Saas zone, which separates it from the overlying garnet-bearing metabasics. These metagabbros are alternatively considered as part of the Combin zone (Forster et al., 2004) or ascribed to the Zermatt-Saas zone, based on petrological and structural data (Pleuger et al., 2007; Groppo et al., 2009). Along its lower contact, the LCU is juxtaposed with a thick sequence of layered metagabbro and antigorite serpentinite of the Zermatt-Saas zone (Forster et al., 2004). Thus, the LCU is enclosed and enveloped by sheared tectonic slices derived from the Zermatt-Saas zone.

4.1.1 Deformation history/style

There is a number of remarkable aspects to the outcrops at Lago di Cignana when it comes to the style of deformation that can be observed. First all deformation that is observed in the Lago di Cignana rocks appears to have taken place during the exhumation. Everywhere, high pressure minerals have been intensely realigned

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during post-P peak deformation and recrystallization. Second it is evident that the rock mass has been boudinaged on all scales - from the kilometer scale, to the outcrop scale, to the microscope scale. In fact it is possible to map boudins within boudins within boudins. There is a history that can be discerned from these overprinting relationships, on the large scale in terms of relative timing, observing how the direction of motion changes, and on the microscope scale, using the combination of petrology, fabric and microstructural analysis, and microstructurally focussed ⁴⁰Ar/³⁹Ar geochronology to determine the temporal variation of pressure- temperature and process along the exhumation path. Boudins can be observed adjacent to the road, and the path leading down to the lake. The shear zones bounding these boudins are complex: for example the observed sense of shear can vary according to which side of a boudin is being observed. In the later part of the history the effect of some Alpine folding can be observed, with extensional boudins being folded and shortened. This is evident on outcrops near the lake shore.

At the map scale, due to this complex deformation history, slivers may be repeated at several different structural levels, attesting to the thrust-related movement involved at different times in this extensional setting. This is most evident when regarding the lithologies of boudins in the Combin Shear zone, which clearly come from a variety of different structural levels. These outcrops suggest what we have previously described (Forster et al., 2004; Lister & Forster, 2009) as a “tectonic shuffle zone” in this location. Alternatively, Beltrando et al. (2014) suggested that the apparent lithological complexity of the Combin Shear zone is largely inherited from pre-Alpine rift-related extensional tectonics, followed by subduction/exhumation-related shearing.

4.1.2 Main lithologies of the Lago di Cignana unit

The UHP unit consists of a basement of glaucophane eclogites with zoisite/clinozoisite + paragonite pseudomorphs after lawsonite, derived from original MOR basalts (cf. Groppo et al., 2009, Figs. 20, 21) and of a metasedimentary cover series. The metasedimentary sequence consists of four main lithologies, i.e.:

1. Glaucophane-garnet calc-micaschists and impure marbles with zoisite/clinozoisite + paragonite pseudomorphs after lawsonite, which include boudins and/or nodules of quartz-bearing garnetite. The eclogite peak assemblage consisted of aragonite (now calcite), coesite (now quartz), phengite, garnet, glaucophane, epidotes, lawsonite (now Zo/Czo + paragonite pseudomorphs), and accessory rutile, zircon, apatite and graphite.

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2. Omphacite-glaucophane-phengite-porphyroblastic garnet manganesiferous quartz-rich micaschists and quartzites with nodules or boudinaged layers of garnetites, where microdiamonds have been discovered (Frezzotti et al., 2011) (see below), and jadeitite nodules. The eclogite peak assemblage consisted of omphacite (now a Na-Ca amphibole-albite symplectite), glaucophane, phengite, coesite (now quartz), garnet, and lawsonite (now a Zo/Czo + paragonite pseudomorph).

3. Omphacite-garnet-glaucophane-chloritoid-lawsonite quartzites and quartz-rich schists with omphacitite nodules. The eclogite peak assemblage consisted of coesite (now quartz), omphacite-aegirinaugite, garnet, glaucophane, zoisite with allanite core, lawsonite (now Zo/Czo + paragonite), chloritoid, carbonate and accessory rutile, apatite, tourmaline, graphite and sulphides.

4. Piemontite-Mn-phengite-porphyroblastic garnet quartzites and quartz-rich micaschists with cm- to dm-wide black nodules. These nodules, which are most likely the sites of the original ferromanganese concretions of the ocean-floor, mainly consist of a mixture of Mn-Fe±Al oxides, garnets with a wide range of compositions of the spessartine-grossular±almandine solid solutions, minor carbonates (calcite, dolomite and rhodochrosite), pink Mn-bearing phengite, piemontite, talc and local tephroite (Mn-olivine), rutile, haematite, tourmaline, apatite and zircon.

4.1.3 The LCU metamorphic evolution

The manganiferous quartzites of LCU (see above, cover lithology 2), first described by Bearth (1967), have been later re-examined by Dal Piaz et al. (1979), but became well known among the students of ultrahigh- pressure metamorphism after the discovery by Reinecke (1991, 1998) of coesite included in a tourmaline crystal.

The first detailed petrographic description of the LCU unit has been given by Reinecke (1991), who studied the manganiferous quartzite. He found that garnet is zoned and that its mineral inclusions and chemical composition record the whole prograde, peak and retrograde evolution: primary quartz is included in the core and outer rim of the garnet, whereas only coesite is found in the inner rim. The peak metamorphic conditions have been inferred from the garnet inner rim, which is a pyrope-rich (up to Prp41) spessartine coexisting with coesite, talc, kyanite, phengite, paragonite, braunite, piemontite, haematite, rutile, dravitic tourmaline, Mg-

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rich ardennite, apatite and zircon, recrystallized at high water activity. A similar history has been reconstructed also for eclogites by van der Klauw et al. (1997). Peak metamorphic conditions have been estimated at 615

± 15 °C and 2.8 ± 1.0 GPa by Reinecke (1991, 1998) from metasediments and Reinecke et al. (1994) from basaltic eclogites. The exhumation history of the unit has been reconstructed on the basis of microstructural analysis of metabasics by van der Klauw et al. (1997). King et al. (2004) showed that garnet from eclogites preserves trace element evidence of prograde discontinuous reactions.

Recently, Angiboust et al. (2009), in a comprehensive study of the Zermatt-Saas unit eclogites through phase equilibrium modelling and Raman Spectroscopy of carbonaceous material, obtained homogeneous peak metamorphic conditions at about 540 ± 20 °C and 2.3 ± 0.1 GPa. Angiboust et al. (2009) suggested that the higher PT conditions estimated from the Lago di Cignana lithologies might be due: 1) to the detachment of a hm- scale tectono-metamorphic unit, later juxtaposed to the Zermatt-Saas ophiolites at 2.3–2.5 GPa; 2) to the lack of quartz inclusions within garnet or omphacite, which prevents the coesite formation; or 3) to minor non- lithostatic overpressure within the Lago di Cignana eclogitic unit, of the same order as those expected in the subduction channel from numerical experiments (c. 10%: Yamato et al., 2007; Raimbourg & Kimura, 2008).

However, Groppo et al. (2009), studying in detail the metamorphic evolution of the meta-ophiolites of LCU and of the adjoining units, concluded that the peak assemblage garnet + omphacite + Na-amphibole + lawsonite + coesite + rutile formed at T~ 600 and P>3.2 GPa (Fig. 19), i.e. just within the diamond stability field (Day, 2012).

These unusually high pressure conditions were later supported by Frezzotti et al. (2011), who discovered microdiamond inclusions in spessartine-rich garnets from the LCU quartzitic sedimentary cover. At temperatures of about 600°C, pressures conditions ≥ 3.2 GPa are also constrained by the graphite-diamond transition curve of Kennedy & Kennedy (1976). It should be noted, however, that the graphite-diamond transition curve has recently been modified and shifted to lower pressures (e.g., Fried & Howard, 2000; Day, 2012) compared to the one proposed by Kennedy & Kennedy (1976). Depending on which diamond-graphite transition curve is selected, minimum P conditions of diamond formation can vary between 2.8 and 3.2 GPa, at 600°C.

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4.1.4 Age of UHP metamorphism

Lu–Hf and Sm–Nd garnet geochronology by Amato et al. (1999) and Lapen et al. (2003) indicated ages of c.

50 and c. 40 Ma for the prograde and peak metamorphism, respectively. Slightly older ages of c. 44 Ma for the UHP peak metamorphism resulted from both in situ U⁄Pb dating of zircon by Rubatto et al. (1998) and

40Ar⁄³⁹Ar high resolution laser dating of phengitic micas by Gouzu et al. (2006). Considering the previously published Sm–Nd and Lu–Hf geochronological data on a UHP eclogite, combined with the evidence that distinct core-to-rim zoning of different REEs in garnet of metamorphic rocks “record different times along a prograde growth history, and consequently may correctly interpret the isochron ages in terms of the garnet growth interval”, Skora et al. (2009) concluded that prograde garnet growth occurred over a ~30 to 40 m.y. interval, i.e. garnet growth started at ~1.1 to 1.4 GPa pressure at ~70 to 80 Ma ago and peak metamorphism occurred at 38 to 40 Ma ago.

4.2 The impure quartzite and the manganiferous garnetite

The impure quartzites including manganiferous garnetites consist of quartz, garnet, minor phengite, partly chloritized green-brownish biotite, piemontite and accessory rutile, opaque ores (most likely Mn-oxides), apatite and a slightly zoned pale green tourmaline. More than one generation of white mica may be recognized based on microstructural position, grain size and interference color. The pale pink manganoan phengite (Reinecke, 1991) crystallized with the uniaxial 3T polytype. The bimodal arrangement of phengite crystals and the local presence of isoclinal fold hinges within the main foliation indicate that the latter was formed through transposition of an older fabric. In the most retrogressed sites, phengite is pseudomorphically replaced by albite. Tourmaline occurs locally as slightly zoned phenoblasts which include piemontite, rutile and opaque ores. Apatite is typically cloudy for the presence of tiny needle-like opaque segregations, elongated parallel to the host crystal c-axis. Rare clinozoisite crystals also occur, usually oblique to the main foliation, with a pale purple radioactive core. Piemontite is often surrounded by a pale yellow green rim of clinozoisite.

The original manganiferous garnetite occurs in quartzite as a few cm-thick discontinuous layers, knots or nodules pink to red-brown in colour that appear to have formed by boudinage of an original layer (Fig. 22).

Locally, nodules appear to be hinges of rootless isoclinal folds. Garnetite consists of an inner portion of massive, close packed, equigranular aggregate of small (c. 50-300 μm across) welded garnets and accessory

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rutile and apatite. Only in places a thin film of quartz is present among garnet granoblasts. The original almost monomineralic garnetite is usually dismembered into fragments with angular shapes separated and cemented by a network of fractures healed with quartz (Fig. 23 left). Locally, this quartz appears to be recrystallized to poikiloblasts up to more than c. 2 cm across, which show a subgrain microstructure whose grain size is similar to that of the surrounding quartzite. Moving further away from the garnetite core, the quartz-garnet ratio increases, but the shape of the nodule is usually still preserved.

Even if the manganiferous garnetite lenses of the LCU are similar to coticule or pseudocoticule from the Belgian type locality (Baijot et al., 2011), they differ as to both the mineral mode and composition. However, it cannot be excluded that the LCU garnetites may have derived from sediments deposited onto the oceanic crust, consisting of radiolarian ooze impregnated with halmyrolized (halmyrolization = a kind of submarine weathering) Mn-rich oxides and/or carbonates together with a clay fraction. The possible derivation of spessartine from a rhodochrosite precursor, suggested by Lamens et al. (1986), has been confirmed by Schreyer et al. (1992).

4.2.1 Garnets of garnetite

In the massive garnetite, where garnets show a granoblastic microstructure and are welded together, each single garnet crystal may be recognized only because the presence of the reddish core rich in inclusions surrounded by a colourless rim. In the more dismembered portions of garnetite single garnets are cemented by intergranoblastic quartz, each garnet consisting of a reddish core usually rich in inclusions of opaque and carbonates, and a colourless inclusion-free or inclusion-poor rim (Fig. 23 left). Locally, garnets may exhibit an atoll-like habit, consisting of a reddish core rich in inclusions with corroded appearance, and a colourless idioblastic rim. Core and rim are separated by a quartz mantle. However, different types of shape and zoning may be observed (Fig. 23 right).

4.2.2 Inclusions in garnet of garnetite

Garnet cores are usually rich in inclusions of opaque, quartz/coesite magnesite, and dolomite (Fig. 23). Mineral inclusions range in size from 40 to 200 µm and show rounded or elongated shapes. Coesite and quartz are commonly present as single crystals, and the usual petrographic features associated with the coesite-quartz transition are lacking (i.e. visible coesite relic; palisade-textured quartz; polygonal granoblastic quartz

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aggregate). Garnet contains also rare microdiamond inclusions. Microdiamonds range from 1 to 20 µm in size, with an average size of 2-6 µm. They are blackish, subhedral to euhedral, and often show well developed octahedral and cubic shapes. Microdiamonds are always associated with carbonates and fluid inclusions (i.e.

coeval). Metamorphic graphite has not been observed. Diamond cubic structure (i.e. C-C bonding of sp³- hybridized C) was confirmed by a sharp band at 1332 (± 2) cm^-1 in the Raman spectra. All diamond bands from a data set of more than 40 spectra show high crystallinity with a Raman band spread of 3.5 - 7.6 cm^-1 FWHM (full with at half peak maximum intensity). Water-rich fluid inclusions (Fig. 23; 2-30 µm in size) associated with microdiamonds contain several daughter crystals, including Mg-calcite/calcite, quartz, rutile, paragonite ± dawsonite ± rhodochrosite ± hydrous Mg-carbonate and sulfate (e.g., dypingite Mg5(CO3)4(OH)2• 5H2O, and pentahydrite MgSO4• 5H2O).

Raman microspectroscopy detected bicarbonate, sulfate, carbonate ions (peaks at 1017, 981, and 1066 cm^-1, respectively), and H4SiO4 monomers and dimers in the aqueous solution (Frezzotti et al., 2011). Raman failed to detect CO2 in gas bubble, constraining the mole fraction of CO2 in the fluid phase to be < 0.026.

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geological field trips 2014 - 6(1.1)itinerary

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Day 1: Punta Rossa unit (Valaisan domain)

STOP 1.1 - Alpine architecture: The Punta Rossa-Hermite-PSB stack (Coord.: Long: 6.870780 E; Lat: 45.695840 N)

This first Stop is located along the main track, in the flat meadow after a calcschist cliff. This stop is aimed at illustrating the overall tectonic architecture of the area. The Petit St. Bernard, Hermite and Punta Rossa units are stacked along south-dipping contacts and the Petit St. Bernard unit occupies the southernmost position (left in Fig. 12). The contact with the underlying Hermite unit is marked by a sharp shear zone dipping to the south, which is locally marked by cargneule, tectonic breccias with angular clasts in a yellow calcite matrix.

An extensional crenulation cleavage is commonly found in the black micaschists and in the Arguerey calcschists of the Hermite unit, in the footwall, indicating that this tectonic contact was characterized by top-to-the-south

Fig. 12 - Landscape view from STOP 1.1

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kinematics (Cannic et al., 1995; Beltrando et al., 2012). The lower boundary of the Hermite unit can be seen at the base of the grassy slope above the grey-greenish outcrops of the Punta Rossa metagranite. From this viewpoint, all the different lithologies of the Hermite unit can be observed. The Versoyen complex, with the characteristic banded appearance related to the green-gray mafic sills and the black schists, can be seen both in the cliffs to the left and on the top of the Aiguille de l’Hermite. The base of the Aiguilles de l’Hermite, as seen from here, consists of Collet des Rousses conglomerates, resting on the overturned limb of a large Frec/F2 fold. Please, note that also this metasedimentary formation is banded and, from a distance, can hardly be distinguished from the Versoyen complex. Collet des Rousses conglomerates crop out also in the foreground, in the cliff directly above the Punta Rossa unit. The Arguerey calcschists, instead, crop out both in the grassy slope in the foreground and on the north-eastern ridge of the Aiguille de l’Hermite. Observations along the Torveraz streambed, in an area not visible from here, indicate that the Arguerey calcschists rest in the core of a tight antiform (Figs. 6 and 7), with axial planar cleavage in the black schists defined by elongated chloritoid and white mica. This fold generation is ascribed to Fctd by Beltrando et al. (2012), which corresponds to F1 in Fügenschuh et al. (1999) and Loprieno et al. (2011). The lower limb of this tight fold, towards the contact with the underlying Punta Rossa unit, is dissected and partly dismembered by minor shear zones (not visible from here). As already mentioned above, the Punta Rossa unit rests directly underneath this late-stage tectonic contact, which will be seen in Stop 1.2. The serpentinites (green) can already be spotted in the rocky ridge in the middle of the picture.

STOP 1.2 - Contact between Punta Rossa and Hermite units (Coord.: Long: 6.862320 E; Lat: 45.697590 N)

Before looking at the internal architecture of the Punta Rossa unit, which is the main topic of this excursion, one more stop is devoted to the understanding of the overall architecture of the area and the contact with the Hermite unit. Stop 1.2 is located along the main track leading towards the Punta Rossa pass, a few tens of meters above the marsh. A view to the south, to the streambed and the cliff lying behind it, allows introducing some of the main lithologies of the Punta Rossa unit, where the Paleozoic basement is overlain by Mesozoic sediments (Fig. 13). The bottom of the valley, floored by pale greenish rocks, consists of metagranitoids, locally preserving the original igneous mineral assemblage, consisting of sub-cm sized quartz, plagioclase and K- feldspar, now largely replaced by aggregates of albite and fine grained white mica. White mica inclusions found

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Fig. 13 - Landscape view from STOP 1.2 (a) and main features of the Punta Rossa grey micaschists in immediate footwall of the contact with the overlying Hermite unit (b and c).